Application to Mount Etna, Italy

Journal of Volcanology and Geothermal Research 251 (2013) 98–111

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An overview of experimental models to understand a complex volcanic instability: Application to Mount Etna, Italy Valerio Acocella a,⁎, Marco Neri b, Gianluca Norini c a b c

Dipartimento di Scienze Geologiche, Università Roma tre, Roma, Italy Istituto Nazionale di Geofisica e Vulcanologia, Sezione di Catania, Catania, Italy Istituto per la Dinamica dei Processi Ambientali, Consiglio Nazionale delle Ricerche, Dalmine, Italy

a r t i c l e

i n f o

Article history: Received 23 April 2011 Accepted 3 June 2012 Available online 13 June 2012 Keywords: Volcano instability Analogue modeling Etna Unbuttressing

a b s t r a c t Volcanic edifices are often unable to support their own load, triggering the instability of their flanks. Many analogue models have been aimed, especially in the last decade, at understanding the processes leading to volcano flank instability; general behaviors were defined and the experimental results were compared to nature. However, available data at well-studied unstable volcanoes may allow a deeper understanding of the specific processes leading to instability, providing insights also at the local scale. Etna (Italy) constitutes a suitable example for such a possibility, because of its well-monitored flank instability, for which different triggering factors have been proposed in the last two decades. Among these factors, recent InSAR data highlight the role played by magmatic intrusions and a weak basement, under a differential unbuttressing at the volcano base. This study considers original and recently published experimental data to test these factors possibly responsible for flank instability, with the final aim to better understand and summarize the conditions leading to flank instability at Etna. In particular, we simulate the following processes: a) the longterm activity of a lithospheric boundary, as the Malta Escarpment, separating the Ionian oceanic lithosphere from the continental Sicilian lithosphere, below the most unstable east flank of the volcano; b) spreading due to a weak basement, with different boundary conditions; c) the pressurization of a magmatic reservoir, as that active during the 1994–2001 inflation period; d) dike emplacement, as observed during the major 2001 and 2002–2003 eruptions. The experimental results suggest that: 1) the long-term activity of a lithospheric tectonic boundary may create a topographic slope which provides a differential buttressing at the volcano base, a preparing factor to drive longer-term (>10 5 years) instability on the east flank of the volcano; 2) volcano spreading (b 104 years) has limited effect on flank instability at Etna; 3) magmatic intrusions (b 101 years), both in the form of Mogi-like sources or dikes, provide the most important conditions to trigger flank instability on the shorter-term. © 2012 Elsevier B.V. All rights reserved.

1. Introduction Volcanic edifices result from the repeated emplacement of magma, so that any edifice with significant height can become unable to support its own load. This lack of support may result in the instability of the volcano flanks. Flank instability is observed at many polygenetic stratovolcanoes, composite volcanoes and shield volcanoes. Destabilization is usually produced by a combination of circumstances and events, rather than a single cause (Voight and Elsworth, 1997). Magma emplacement is the most common triggering factor (Voight et al., 1981; Delaney et al., 1998; Belousov et al., 1999; Elsworth and Day, 1999; Richards and Villeneuve, 2001; Tibaldi, 2001; Acocella et al., 2006). Other common triggers are fault activity (Hall et al., 1999; Walter et al., 2005; Norini et al., 2008), earthquake shakes (Ando, 1979; Acocella et al., 2003), a ⁎ Corresponding author. E-mail address: [email protected] (V. Acocella). 0377-0273/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2012.06.003

weak basement (Borgia et al., 1992; Morgan et al., 2000; Norini et al., 2010) and hydrothermal alteration (Lopez and Williams, 1993; Reid et al., 2001; Behncke et al., 2008). Modelling allows understanding mechanisms related to flank instability. Analogue models have been widely used to simulate and understand general processes of volcano instability as a function of the cone height and the brittle-ductile ratio of the substratum (Merle and Borgia, 1996; van Wyk de Vries and Merle, 1996), of a basal ductile horizon (Merle and Lenat, 2003), cryptodomes (Donnadieu and Merle, 1998, 2001; Acocella, 2005) or of a basal fault (van Wyk de Vries and Merle, 1998; Lagmay et al., 2000; Vidal and Merle, 2000; Merle et al., 2001; Acocella, 2005; Norini and Lagmay, 2005; Norini et al., 2008). Here we use analogue models to consider flank instability in a specific, relatively well-known volcano, providing indications on the driving mechanisms at the local scale. We expect that a better definition of the instability processes occurring in a known volcano may

V. Acocella et al. / Journal of Volcanology and Geothermal Research 251 (2013) 98–111

help in understanding general processes of flank instability also at less studied or poorly monitored volcanoes. In this frame, Etna (Italy) provides a suitable case. In fact, on the one side, Etna is characterized by a well-monitored flank instability and a definition of the kinematic features of its flanks (e.g. Solaro et al., 2010, and references therein). On the other side, these kinematic features are still not properly placed in a wider context, as different triggering factors of flank instability have been proposed. A weak, clayish basement has been first interpreted as the main reason for flank instability (Borgia et al., 1992; Tibaldi and Groppelli, 2002; Rust et al., 2005). InSAR data suggest that the differential buttressing conditions at the volcano base may also play an important role (Froger et al., 2001). In a similar fashion, the role of the instability of the nearby continental margin has been inferred to drive the long-term seaward slide of the eastern flank of the volcano (Chiocci et al., 2011). After the major 2001 and 2002–2003 eruptions, the role of shallow magma emplacement, mainly in the form of dikes, on volcano instability was highlighted (Acocella et al., 2003; Puglisi et al., 2008; Neri et al., 2009; Ruch et al., 2010). Modeling of InSAR and GPS data has shown that magma pressurization from deeper sources may also induce flank instability (Bonforte and Puglisi, 2006; Palano et al., 2008). In addition, field and seismicity data highlight the role of regional tectonics, related to the development of the Malta Escarpment (ME), on the instability of the flanks of the volcano (Monaco et al., 2005). In this frame, a recent overview of InSAR data from the 1994–2008 period shows how the flanks of Etna undergo a complex mechanism of instability, resulting from the interplay of at least 3 processes, under conditions of differential unbuttressing (or confinement conditions) at the base of the volcano. These processes are: the load of the volcano, inflation due to magma accumulation at depth and dike emplacement (Fig. 1; Solaro et al., 2010). Given this large variability in the processes possibly associated to flank instability, it becomes essential to try to provide a hierarchy of importance of these factors. This study describes original (experimental sets 1 and 2, see below) and reviews recently published (experimental sets 3 and 4, see below, published in Norini and Acocella, 2011) experimental data. The goal is to test the feasibility, the role and the implications related to the processes possibly

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responsible for flank instability, providing a summary for these conditions. To this aim, we present four main sets of analogue experiments, capturing the possible role of long-term tectonic activity, differential buttressing conditions, a weak basement, the pressurization of a magmatic reservoir, dike emplacement. The experimental results confirm the importance of at least four of these processes on flank instability at Etna. 2. Structure of Etna and of the unstable flanks Etna is located at the northern edge of the Sicilian foredeep, between the Appennine–Maghrebian Chain (AMC) to the north, and the Hyblean Foreland (Hy) to the south. Etna is also located along the NNW–SSE trending ME, a lithospheric boundary separating the Ionian oceanic crust from the Sicilian continental one, and lies on Pre-Quaternary foredeep and Plio-Pleistocene clayish deposits (Fig. 2, inset; Barberi et al., 1974; Monaco et al., 1997; Gvirtzman and Nur, 1999; Chiocci et al., 2011). The NNW–SSE trending rightlateral transtensional Timpe Fault System (TFS) represents the onshore part of the ME on the E flank of the volcano (Fig. 2) (Monaco et al., 1997, 2005; Corsaro et al., 2002; Azzaro et al., 2004; Siniscalchi et al., 2012). To the north, the TFS meets the sinistral transtensive NNE–SSW Messina Faults System (MFS), which marks the NEtrending Ionian coastline of Sicily from Etna towards the Messina Strait (Lanzafame et al., 1996, 1997, and references therein). The structure of the volcanic edifice is dominated by three main “rift zones”, radiating from the summit (Rittmann, 1973): the NE Rift, the S Rift and the less developed W Rift (Fig. 2); these are usually fed by the central conduit, through the lateral propagation of radial or subradial dikes, rather than through vertically propagating dikes from an underlying shallow magma chamber (Sharp et al., 1980; Sanderson, 1982; McGuire and Pullen, 1989; Bousquet and Lanzafame, 2001; Patanè et al., 2003). Etna is characterized by the outward displacement of its flanks, most notable on the faster eastern and southern sides; here average rates range between ~0.5 and ~5 cm/yr (Kieffer, 1985; Borgia et al., 1992; Lo Giudice and Rasà, 1992; Lundgren and Rosen, 2003; Mattia et al., 2007; Solaro et al., 2010; Bonforte et al., 2011). The unstable

Fig. 1. Main collapse processes responsible for the 1992–2008 deformation observed on Mt. Etna (Solaro et al., 2010). a) Effect of the load of the volcano; b) effect of magma accumulation beneath the summit; c) effect of dike emplacement during 2001 and 2002.

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Fig. 2. Simplified structure of Mt. Etna (a), reporting the main rifts, the main fault zones and the area affected by major flank instability (grey); black arrows indicate the directions of movement of the slide blocks within that area. TFS = Timpe Fault System, RFS = Ragalna Fault System, PFS = Pernicana Fault System, MSF = Messina Fault System. Inset (b): regional tectonic setting; AMC = Apennine–Maghrebian Chain; CGF = Catania-Gela Foredeep; HF = Hyblean Foreland; ME = Malta Escarpment; CoF = Compressional Front; EF = Extensional Front.

area is confined to the N by the E–W trending transtensive left-lateral Pernicana Fault System (PFS) and to the SW by the N–S trending transtensive right-lateral Ragalna Fault System (RFS) (Fig. 2) (Rust and Neri, 1996; Tibaldi and Groppelli, 2002; Acocella and Neri, 2005; Rust et al., 2005; Neri et al., 2007; Siniscalchi et al., 2010, 2012). In between these faults, several fault systems (mostly coinciding with the above mentioned Timpe Fault System), with an overall N–S to NW–SE trend and predominant transtensive right-lateral kinematics, are active; these usually bound different sectors or, in some cases, rigid blocks, moving with differential kinematics; the overall moving direction of these blocks on the east flank is ESE (Fig. 2) (Borgia et al., 2000; Froger et al., 2001; Acocella et al., 2003; Lundgren et al., 2004; Neri et al., 2009; Solaro et al., 2010). A close relationship between flank deformation and eruptive activity at Etna has been recognized (e.g. Acocella et al., 2003; Burton et al., 2005; Neri et al., 2005, 2009; Bonaccorso et al., 2006; Bonforte et al., 2007, 2008, 2009; Puglisi et al., 2008). In the last two decades, the 2001 and, mostly, 2002–2003 flank eruptions, both fed by a vertically-propagating N–S trending dike, were accompanied by destructive seismic activity, extensive surface fracturing, and accelerated seaward sliding on the E flank, allowing a better understanding of the gravitational and magmatic processes (Acocella and Neri, 2003; Acocella et al., 2003; Behncke and Neri, 2003; Bonforte et al., 2004, 2009; Neri et al., 2004, 2005; Aloisi et al., 2006). This correlation between volcanic and instability events shows the importance of

defining further the conditions leading to flank instability, to better understand and forecast major eruptions at Mount Etna. 3. Experimental apparatus and scaling 3.1. Boundary conditions and apparati Based on the study of Solaro et al. (2010), the role of the following processes has been tested and studied, in four main experimental sets, to study the complex instability at Mt. Etna. 1) The long-term activity (>10 5 years) of the ME and MSF below the E flank of the volcano; these fault systems are associated with the development of the lithospheric trenstensive Malta Escarpment (e.g. Gvirtzman and Nur, 1999; Argnani and Bonazzi, 2005; Chiocci et al., 2011; Siniscalchi et al., 2012), and an overall E–W extension (Fig. 3a; Lanzafame et al., 1997; Monaco et al., 1997). Such an activity marks the transition from a thicker continental crust, to the west, to a thinner oceanic crust, to the east. An important outcome is the development of a topographic gradient in the volcano basement (mean slope angle of 5°–10°, over a distance of ~50 km), marking the transition from the mean altitude of the Apennine-Maghrebian Chain (1 km a.s.l., to the N and W) to the bottom of the Ionian oceanic abyssal plain (−2 km b.s.l., to the E) (e.g. Froger et al., 2001). The effect of the long-term activity of


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to the widespread presence of the Miocene–Pliocene clayish and sandish flysch deposits below the entire edifice, with a thickness of a few kilometers (Lentini, 1982). To this aim, we present an original experimental set (set 2), simulating the effect of such a weaker basement. In this set, a silicone basement is placed at the base of a cone under various types of differential confinement, to simulate differential buttressing conditions at the weak volcano base. 3) The emplacement of magma within the volcano basement has been simulated considering the pressurization of deep intrusions, as the ones modeled as Mogi-like reservoirs by Palano et al. (2008), active between 1993–2001 and occurring on the shortterm (10 1-2 years) (Fig. 3c). The pressurization of the magmatic system has been usually simulated injecting higher or lower viscosity silicone putty through a nozzle (with a diameter of 1 cm) at the base of the model. This set of experiments is described in more detail in Norini and Acocella (2011), and only a summary is presented here. This experimental set (set 3) is thus characterized by the simulation of a pressurized magma reservoir within a volcanic cone above an asymmetric (partly inclined) basement, reproducing the topographic/bathymetric gradient at the volcano base. 4) The emplacement of feeder dikes in the upper part of the volcano, acting on the short-term, as during the 2001 and the 2002–2003 eruptions (Acocella et al., 2003; Puglisi et al., 2008; Neri et al., 2009), was simulated injecting low-viscosity vegetable oil with a peristaltic pump, through a 5 mm diameter nozzle, at a known rate (Fig. 3c). The intrusions were stopped when the oil reached the surface of the model. This experimental set (set 4) is thus characterized by the simulation of dikes emplaced within a volcanic cone above an asymmetric basement. This set of experiments is described in more detail in Norini and Acocella (2011) and here only a summary is presented. 3.2. Scaling and materials

Fig. 3. Apparati and boundary conditions used in the four sets of experiments. a) Velocity Discontinuity (VD) at the base of a granular cone, responsible for the creation of localized extension. b) Weak basement (due to a layer of silicone) imposed at the base of a granular cone. Case 2a is totally unconfined, cases 2b and 2c are partially confined. c) Intrusion of vegetable oil and silicone to simulate the emplacement of dikes and the pressurization of shallow magmatic sources, respectively.

ME and MSF on flank instability is investigated through an original experimental set (set 1), characterized by the sliding of a basal discontinuity, with a kinked shape, below a cone, creating a set of normal/transtensive faults responsible for the development of the topographic gradient. 2) The possible role of a weak basement below the volcanic pile, acting on the mean-term (b104 years), based on suggestions from previous studies and on the onset of activity of the Pernicana Fault System on the present volcanic edifice (Mongibello) (Fig. 3b; e.g. Borgia et al., 1992; Tibaldi and Groppelli, 2002; Solaro et al., 2010). Considering the local geology, this basement may be related to the presence of Plio-Pleistocene clays, immediately at the base of the volcanic pile, on the eastern and southern base of the volcano (Aa.Vv., 1979), and

Models have to be geometrically, kinematically and dynamically scaled, following the principles discussed by Ramberg (1981). We chose a length ratio between model and nature z⁎ = 2 ×10 − 5 (1 cm in the model corresponds to approx. 500 m in nature; Table 1). The densities of natural volcanic rocks (2000–2700 kg m − 3) and of the commercially available experimental materials (900–1800 kg m − 3) impose a density ratio ρ* ~ 0.5. Since the models were run at 1 g, the gravity ratio g* = 1. These ratios imply that the stress ratio between model and nature is σ* = ρ*g*z* ~ 10 − 5 (Table 1). We assumed a Mohr–Coulomb failure criterion for the rocks of the volcanic edifice, with an angle of internal friction ϕ = 35° and a mean cohesion c between 10 6 and 10 7 Pa. Cohesion, having the dimensions of a stress, must be scaled at ~10 − 5 in the experiments; this requires the use Table 1 Values and scale factors of the main parameters used for scaling. 1: granular materials (mixture of sand and flour, and mixture of silica sand and crashed silica powder); 2: vegetable oil; 3: higher-viscosity silicone; 4: lower-viscosity silicone.

1

2

3 4

Parameter

Nature

Model

Scale factor

Scale

Length (m) Density (kg/m3) – Cohesion (Pa) Velocity of intrusion (m/s) Viscosity of (Pa s) Viscosity of intruding fluid (Pa s) Viscosity of weak basement (Pa s)

3400

10− 2

Length ratio z*

~ 10− 5

2600

~ 1400

Density ratio ρ*

~ 0.5

– 106–107

– 30–200

Stress ratio σ* Cohesion ratio c*

~ 10− 5 ~ 10− 5

0.1

~ 10- 4

Velocity ratio v*

~ 10− 3

104

~ 10- 2

Viscosity ratio η*

~ 10− 6

1015

9 × 105

Viscosity ratio η*

10− 10

1015

1 × 104

Viscosity ratio η*

10− 12

102


of material with c between 10 and 100 Pa to simulate the volcanic edifice. The simulation of a material whose rheology is time-dependant, such as magma or weak basements, has to take into account for the viscosities and the related strain rates, velocities, and time-scales. These parameters are related to each other. Definition of the stress ratio σ* between model and nature allows evaluating, if the viscosity ratio η* in model and nature is known, the corresponding strain rate ratio ε*, according to ε* ~ σ*/2η*. The strain rate ratio ε*, being measured in s − 1, is the inverse of the time ratio ε* = 1/t* and, being multiplied for the length ratio z*, also allows defining the velocity ratio v* = ε*z* (Merle and Vendeville, 1995). We use granular materials to simulate the volcano and the upper portion of its basement (b7 km). In experimental sets 1) and 2), the cone is made up of a mixture of well sorted, round grain, dry quartzose sand (87% in volume) and flour (13% in volume). These proportions give the mixture an angle of internal friction ϕ~35° and cohesion c= 100–200 Pa. Similar mixtures have been previously used (Donnadieu and Merle, 1998; Vidal and Merle, 2000; Merle et al., 2001; Merle and Lenat, 2003). The mixture was released from a funnel at fixed height, building the cone with a natural repose angle. In experimental sets 3) and 4), the granular material consisted of a mixture of silica sand and crushed silica powder. The sand is non-cohesive; it has smooth and rounded grains, with an angle of internal friction Φ ~25° and grain size of 100–200 μm. The crushed silica powder has angular grains with finer size (b50 μm), cohesion c ~300 Pa, angle of internal friction Φ ~40°, and tensile strength of ~100 Pa. A mixture of 90% silica sand and 10% crushed silica powder has cohesion c~30 Pa and tensile strength of ~10 Pa, so that the ratio between shear strength and tensile strength is 3:1. Such a tensile strength ensures propagation of the low-viscosity vegetable oil used to simulate dikes (see below) through hydrofractures (Galland et al., 2006). The possibility to simulate dikes justifies replacing the sand-flour mixture with the crushed sand. The two granular materials we used have slightly different cohesion (30 and 100–200 Pa) which is not expected to influence the deformation pattern in our experiments. The slope of the granular cone in the various experiments varied between 15° and 35°; the deformation pattern was independent of the slope, provided that realistic (scaled) deformative conditions were imposed (Norini and Acocella, 2011). Newtonian silicone putty with different viscosity is commonly used to model weaker deposits (i.e. clays, evaporites) and deep pressurized magma intrusions, consistently with previous experiments (e.g. Merle and Borgia, 1996; Donnadieu and Merle, 1998; Acocella et al., 2001). The overall rheology of the silicone has been discussed in Weijermars et al. (1993) and Hailermariam and Mulugeta (1998). We have used two different types of silicone to simulate the behavior of the weak basement and of the magma (Table 1). A lower-viscosity silicone (~10 4 Pa s) is used to simulate a weak basement (experimental set 2). Given the duration of basal spreading in the experiments (7 × 10 3 s) and in nature (~3.1 × 10 10 s), the time ratio between model and nature t* = 2 × 10 − 7, corresponding to a strain rate ratio ε* = 1/t* = 5 × 10 − 6. This, given the stress ratio σ* = 10 − 5, corresponds to a viscosity ratio η* = 2 × 10 − 12, or to a natural viscosity of the basement η ~ 10 15 Pa s, in agreement with recent estimates (Palano et al., 2009). A higher-viscosity silicone (~9 × 10 5 Pa s) is used for magma intrusions. To simulate the pressurization of a magmatic reservoir, the injection velocity of silicone varied between 1.3 × 10 - 6 and 3 × 10 − 5 m/s, with a duration of a few hours: in particular, we simulated a viscous magma (10 11–10 14 Pa s) to reproduce a pressurized ~ 7 km deep reservoir below Etna, over a time span of 10 1 to 10 2 years. The simulated 10 11–10 14 Pa s viscosity of the pressurized reservoir appears relatively large. Nevertheless, as the aim of our experiments is to study the finite deformation due to the pressurization of a magma analogue, in a first approximation, we neglect any possible discrepancy due to the viscosity of the source and focus on the effect of this pressurization only.

Finally, very low-viscosity molten vegetable oil (~10 − 2 Pa s) simulates shallow dike emplacement (experimental set 4; see Galland et al., 2006). The oil is solid at room temperature and melts to nearlyNewtonian low-viscosity fluid with density of ~0.9 g/cm 3 when the temperature rises above 28 °C. The oil is intruded in the model with a peristaltic pump, with a rate of intrusion of 2 × 10 − 2–6 × 10− 2 m/s. When the oil propagates and intrudes the model, the velocity lowers to about 10 - 4 m/s (Table 1). The oil is able to intrude when it is at 35 °C, with a viscosity of ~2 × 10− 2 Pa s. Both the length ratio z* and the stress ratio σ* are in the order of 10− 5. To scale velocity, we assume low viscosity intrusions in nature propagate at ~0.1 m/s (Spence and Turcotte, 1985; Battaglia and Bachelery, 2003; Roman et al., 2004). The viscosity of the molten vegetable oil we used in the models is 10− 2 Pa s and corresponds to magma viscosity of 104 Pa s in nature, which is in the range of the basaltic magma of Etna (Spera, 2000; Harris and Allen, 2008). The chemical incompatibility between the silica grains and the vegetable oil, as well as the fine grain size of the crushed silica powder, prevents percolation of the fluid in the granular material. The intrusion of the vegetable oil simulated the emplacement of dikes in the Etna edifice over a time span of a few days, consistently with observations (e.g. Neri et al., 2005). 4. Experimental Results The analogue models are divided into 4 main sets, depending upon the imposed conditions. The experimental results have been studied observing both the surface (incremental deformation) and the internal part (final deformation) of the models. All the experiments were photographed to record the structures developed during progressive deformation. In order to capture the more detailed (mmscale) deformation of sets 3 and 4, we analyzed the surface data of these models with a laser scanner in a Geographic Information System (GIS) (Norini and Lagmay, 2005). Incremental and final elevation changes were measured through digital elevation models (DEMs) from laser scanner data referred to a fixed reference (the location of the laser scanner), to depict deformation areas with a resolution of 10 2 μm. Completed experiments were wet and then sliced to reveal the structures within the cone and basement in cross-sections. In the following sections, the main features of each experimental set are described. 4.1. Set 1: the activity of a lithospheric boundary This modeling set consists of 5 experiments simulating the activity of the lithospheric boundary which includes ME and MFS, below the E flank of Etna. We consider the extensional structures along the coastline of E Sicily and below the E flank of Etna having a NNE–SSW to NE–SW trend to the N (MSF), and NNW–SSE to the S (ME; Monaco et al., 1997; Nicolich et al., 2000; Argnani and Bonazzi, 2005; Monaco et al., 2005, 2008). In the experiments, we impose a basal velocity discontinuity (VD) with a change in direction (from NNE–SSW to NNW–SSE) below the flank of a cone (Fig. 4a). The VD is given by an outward sliding basal plate, simulating the localized activity of an extensional fault zone, accordingly with previous experiments (e.g. Mauduit and Dauteuil, 1996; Acocella et al., 1999; Acocella, 2005). Conversely from Acocella, 2005, here the basal discontinuity has a kinked, not rectilinear, shape (see VD in Fig. 4a). We consider a resulting overall extensional motion along an E–W direction for ME and MFS, consistently with available data (Lanzafame et al., 1997; Monaco et al., 1997, 2005; Argnani and Bonazzi, 2005). Two representative experiments are here shown, depending on the distance D between the cone summit and the nearest edge of the basal discontinuity. In experimental set 1a, a 8 cm high cone (H) was resting on a 4 cm thick basement, above the VD (Fig. 3a). The minimum distance between the cone summit and the VD was D = 5 cm. The trace of the


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Fig. 4. Representative experiments from Set 1, characterized by different values of the VD-summit distance D (see below). a), b) and c) Map views of a cone undergoing progressive basal failure (trace of VD shown in a)) at 0, 7 and 25 min, respectively. The minimum horizontal distance between the VD and the cone summit is D = 5 cm. X and Y are synthetic and antithetic faults, respectively. d), e) Map views of a cone undergoing progressive basal failure (trace of VD shown in d)) at 0 and 25 min, respectively. The minimum horizontal distance between the VD and the cone summit is D = 10 cm. f) Oblique view showing a detail of e) (view from top of e)). White arrows in a) and d) indicate direction of extension of the VD.

VD is represented by the white dashed line in the undeformed stage (Fig. 4a). The mobile plate (velocity = 5 cm/h) induces a fault propagating upward from the VD (t = 7′; Fig. 4b). The ~ 50° dipping fault (X in Fig. 4b) reaches the summit with an arcuate shape, corresponding to the intersection with the topography. Its kinematics varies from dipslip (cone summit) to left-lateral and right-lateral (cone base; Fig. 4b), similarly to Acocella (2005). However, conversely to Acocella (2005), here the kinked configuration of the VD allows reproducing an asymmetric (with regard to the extension direction; white arrow in Fig. 4a) collapsed area delimited by the fault. The capability of the fault to reach the cone summit depends on its dip, the initial distance of the VD from the summit (D) and the cone height (H); in experiment 1a, H/D = 1.6; this value falls in the range (1.3 b H/D b 1.7) where the summit of a cone and the instability of a flank may be affected by basal faulting (Acocella, 2005). The enlargement of the fault affected area towards the summit results from two factors: a) the higher thickness of the cone with regard to the basement, (e.g. Mauduit and Dauteuil, 1996); b) the gravitational stresses enhancing fracturing on the cone summit (Acocella, 2005). After 7 min, the fault scarp enlarges without structural variations (Fig. 4c). Experimental set 1a has shown that basal faulting may affect the summit and therefore the stability of an entire flank of a cone, with an asymmetric pattern with regard to the extension direction. However, since H/D = 0.35 at Etna, it is unlikely that basal faulting may affect the cone summit. This is confirmed by experiment 1b, which is characterized by a more realistic H/D = 0.8, given by the height of the cone H = 8 cm and the distance between the VD and the cone summit D = 10 cm (Fig. 4d). Even though this ratio still does not match that Etna, it is interesting to see how, despite the overall similar evolution of the deformation, in this case the portion of the

cone affected by faulting is significantly reduced. This is shown in the final stage of deformation of the experiment (Fig. 4e), where the faulted area is limited to the lower slope of the cone. The gravitational effect of the lower faulted part of the cone develops a very localized area of instability, which does not reach the mean portion of the slope. In addition, the base of the cone is characterized by the development of a few cm deep graben, with a steep slope (Fig. 4f) and strike parallel to that of the underlying VD. In experiment 1b, the basal plate moved outward of 2.5 cm, simulating a depression with a vertical displacement of a few kilometers, consistently with that accumulated along the northern part of the ME over a time span of a few Ma (Argnani and Bonazzi, 2005, and references therein). This experimental set, even though suggesting a limited effect of faulting on the deformation of the cone for Etna, shows how the activity of a lithospheric boundary with an extensional component may create a significant depression, bordered by a steep slope on the edge of a volcanic edifice. 4.2. Set 2: the role of a weak basement This modelling set consists of 5 experiments testing the role of a weak layer at the base of a cone. Three representative experiments are here considered: the spreading of a cone with circular base on a totally unconfined weak basement (2a); the spreading of a cone with circular base on a partially unconfined weak basement (2b); the spreading of a cone with elliptic base on a partially unconfined weak basement (2c). The conditions of experiment 2a are similar to what previously presented (e.g. Merle and Borgia, 1996; Wooller et al., 2004). Here the cone is 8 cm high, with a basal diameter of 18 cm, giving its

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Fig. 5. Representative experiment from Set 2, with totally unconfined conditions (case 2a in Fig. 3b). a), b), d) Map view of the experiment at 0, 15 and 110 min, respectively; c) detail of d). e) Map view of the silicone base after the removal of the granular material, at the end of the experiment; the tangential basal anticline is evident. f) Variation of the height (grey) and width (black) of the cone with time.

slope an initial dip of ~40°; the cone is resting over a 1 cm thick layer of low viscosity silicone (Fig. 3b). This thickness is justified by the fact that any sliding surface(s) at Etna is not shallow, and may lie well below the volcanic pile (e.g. Ruch et al., 2010; Siniscalchi et al., 2012, and references therein). The thickness of the brittle layer overlying the silicone varied between 0 and 1 cm, accordingly with the large variability in the thickness of the volcanic deposits at Etna (Neri and Rossi, 2002). The undeformed map view of the model is shown in Fig. 5a. Soon the cone starts to deform, its load over the weaker basement inducing a decrease in height and increase in diameter. At approximately 15 min, several normal faults start to develop on the outer periphery of the cone, to accommodate its enlargement (Fig. 5b): the faults intersect each other at ~45°, converging towards the outer side of the cone and bordering grabens whose axis is radial to the cone. A detail of these structures, at the final stage of deformation, is anticipated in Fig. 5c. At approximately 30 min, the basement below the outer base of the cone starts to deform and a tangential fold starts to appear within the silicone. The final stage of the deformation of the model (110 min) in map view is shown in Fig. 5d. Here the cone has considerably decreased its height and increased its diameter. At its base, the tangential fold in the silicone is well developed and fully visible when the granular cone is removed (Fig. 5e). The overall exponential decrease in the cone height (grey data), as well as its exponential increase in diameter (black data), are illustrated in Fig. 5f.; the approximate times for the onset of the formation of the radial grabens, as well as of the subsequent basal folds, are indicated. Experimental set 2b simulates the spreading of a cone with a circular base on a partially unconfined weak basement, so that three sides of the basal silicone are laterally confined and the silicone flows laterally on the fourth side only. These conditions have been

chosen as they may better simulate the major eastward slip of Etna (Froger et al., 2001; Acocella et al., 2003). The undeformed map view of the model is shown in Fig. 6a. Similarly to experimental set 1a, soon the cone starts to deform, decreasing its height and increasing its diameter. At t = 10 min, the summit of the cone shows an asymmetric graben, whose axis is perpendicular to the spreading direction (Fig. 6b). The normal fault bordering the graben on the unconfined side is steeper than that bordering the graben on the confined side. At t = 12 min, two strike-slip faults (dextral and sinistral) radiate from the points of intersection of the two normal faults at the base of the graben, towards the extremities of the unconfined side (black arrows in Fig. 6b), accommodating the lateral movement of the cone base. The portion of cone towards the confined side does not show any cm-scale deformation, while that towards the unconfined side deforms significantly. On this mobile portion, at t = 20 min, radial graben start to appear (white arrows in Fig. 6b), accommodating the spreading. At the final stage of deformation, after 120 min, the mobile part of the cone shows at least 4 major radial graben intersecting each other (Fig. 6c). The cone base has an elongation towards the unconfined area. The major axis of the cone increases exponentially its diameter, whereas the summit height decreases exponentially (Fig. 6d). However, because of the higher confinement, both the summit subsidence and the base enlargement are reduced with regard to experiment 2a. Experimental set 2b has shown how a cone may spread elongating parallel to the spreading direction. At Etna, the volcanic edifice has a N–S trending elongation, perpendicular to the spreading direction (approximately E–W). To better evaluate the relationships between volcano elongation and spreading, we have considered the spreading of a cone with an elliptic base on a partially unconfined weak basement, where the cone elongation is perpendicular to the spreading


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Fig. 6. Representative experiment from Set 2, with partially unconfined conditions (case 2b in Fig. 3b). a), b), c) Map view of the experiment at 0, 10 and 120 min, respectively; black arrows = trace of incipient strike-slip fault; white arrows = trace of incipient radial graben. d) Variation of the height (grey) and width (black) of the cone with time.

direction (experiment 2c). The elongated cone is 7.5 cm high, with a major axis M at the base of 22.3 cm and a minor axis m of 18.5 cm, with a basal eccentricity m/M ~ 0.8, similar to that of Etna (Fig. 2). All the remaining conditions, including the partial basal confinement, are the same as those of experiment 2b. The undeformed map view of

the model is shown in Fig. 7a. The overall evolution of the experiment resembles, in general, that of experiment 2b. The onset of the formation of the strike-slip faults at the cone base and of the radial graben on the cone are observed at 15 and 17 min, respectively (Fig. 7b, taken at 17 min). The final stage of deformation, at t = 120 min,

Fig. 7. Representative experiment from Set 2, with partially unconfined conditions and elongated cone (case 2c in Fig. 3b). a), b), c) Map view of the experiment at 0, 10 and 120 min, respectively; black arrows = trace of incipient strike-slip fault; white arrows = trace of incipient radial graben.

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inclined (~10°) on the opposite side (Figs. 3c and 8a). The intrusions correspond to less than a very few km 3 in volume and occur within 10 1–2 years, to simulate, for example, the intrusions responsible for the inflation of the volcano between 1993 and 2001, modelled with the pressurization of Mogi-like sources by Palano et al. (2008). In the representative experiment, the high-viscosity silicone putty is intruded through a vertical conduit, 1 cm in diameter, at the base of the brittle basement exactly below the cone summit (Fig. 8a). The thickness of the basement is 7 cm, and the height of the cone is 3 cm. The intruded volume is 2 cm 3, corresponding to 2 km 3 and approximately to 10 1–2 years of activity at Etna (Allard et al., 2006; Neri et al., 2009). The symmetric intrusion induces asymmetric flank instability, more evident towards the basement slope (Fig. 8a). At earlier stages of deformation (b1 cm 3 of intrusion), the cone surface exhibits incipient extensional fractures at the summit, striking perpendicular to the direction of the basement slope. After 1 cm 3 of silicone intrusion, internal deformation generates folding at the cone base, further opening the summit extensional fracture, and developing strike-slip faults on the cone flank. Progressive intrusion develops an incipient fractured and faulted collapsing sector, with the bisector towards the basement slope (Fig. 8a). The analysis of the surface displacement map shows strong asymmetry of the deformation during intrusion (Fig. 8b). The flank opposite to the basement slope exhibits no evident displacement. The cone summit subsides up to − 1.9 mm, whereas the flank towards the basement slope deforms up to 2 mm (Fig. 8b).

4.4. Set 4: dike emplacement

Fig. 8. Deformation of set 3 experimental models subjected to high-viscosity silicone intrusion. a) Map view of the model after 2 cm3 of silicone intrusion. The cone shows summit extensional fracture, strike-slip faults on the flanks, and folding at its base. These structures comprise an incipient collapse area, with direction of instability towards the basement slope. b) Displacement map showing the net (vertical and horizontal) surface deformation of the model induced by the silicone intrusion, depicted as shades of red and blue. The net deformation resulted from both vertical and horizontal movements of the cone flank.

shows similarities with the previous experiment, except for the more evident development of structures striking parallel to the original cone elongation, towards the unconfined side of the model (Fig. 7c). The final base of the model is elongated parallel to the spreading direction, that is perpendicular to its original elongation. 4.3. Set 3: intrusion of magma in the basement This modelling set consists of 12 experiments and is described in more detail in Norini and Acocella (2011). Here, for the benefits of comparison with the new results, we summarize the evolution of a significant experiment of intrusion within an asymmetric basement. The latter may result from the activity of the lithospheric discontinuity described in set 1 (Fig. 3c). In the experiment, the basement of the cone is horizontal on a side of the injecting nozzle, and slightly

This set consists of 8 experiments, and is described in detail in Norini and Acocella (2011): here, for the benefits of comparison with the new results, we summarize the previously described results. The molten vegetable oil is intruded in the model, forming dikes propagating upward, from the basement to the cone (Fig. 3c). The aim of this set is to analyze the geometry of the dikes and to study the deformation induced by their emplacement. Consistently with results from experimental set 1, a topographic gradient has been imposed to the basement of the cone (Fig. 3c). The basement, similarly to set 3, is horizontal on a side of the injecting nozzle, and slightly inclined (~10°) on the opposite side (Figs. 3c and 9). The intrusion of the vegetable oil occurs from a 5 mm diameter circular orifice centred beneath the cone summit; injection continues until the oil reaches the surface, usually within 60 s (Fig. 9a). No constrain on the direction has been imposed to the intruding oil, and the resulting dikes propagate in the brittle granular material through hydraulic fracturing, following the stress distribution in the model. In the experiments, the thickness of the basement is 5–7 cm, and the height of the cone is 2–5 cm, in agreement with available data showing that dike emplacement affects the volcano basement (e.g. Bonaccorso et al., 2002; Currenti et al., 2010). The results invariably confirm that the injection of the molten oil generates a single dike, propagating from the model base to its surface (Fig. 9a; Galland et al., 2006). The analogue cones are characterized by extension trending perpendicular to the slope direction during intrusion. The flank of the cone on the basement slope shows the most significant movement, with an elevation increase of approximately 1.2 mm; the summit of the cone and the flank opposite to the basement slope show an elevation decrease of approximately −0.8 mm (Fig. 9b) (Norini and Acocella, 2011). These changes have been detected combining vertical and horizontal movements of the cone flanks and summit (see Norini and Acocella, 2011). Considering the model scale and frequency of dike injection at Etna, the deformation rate observed in this set corresponds to 100–10 1 cm/yr in nature (Norini and Acocella, 2011). Serial cross-sections of the models parallel to the basement slope show a single subvertical dike, up to 5–10 cm long and usually


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Fig. 9. Deformation of set 4 experimental models subjected to intrusion of low-viscosity vegetable oil. a) Map view of the model with the extruded oil flowing on the cone flank. b) Measured net (vertical plus horizontal) deformation of the cone induced by the oil intrusion, depicted as shades of red and blue. The trace of the cross-section is shown. c) Internal deformation of set 4 experimental models subjected to low-viscosity vegetable oil intrusion. The cross-section of the model shows dike and sill propagating into the model.

2–3 mm thick, as for example reported in Fig. 9c. The direction of the dike was always perpendicular to the basement slope. 5. Discussion 5.1. Interpretation of the experimental sets Experimental set 1 evaluated the possible longer-term (>105 years) effects of the activity of ME and MFS on flank instability at Etna. Interestingly, conversely to what observed in previous experiments (Acocella, 2005), here the asymmetric shape of the collapsed area with regard to the extension direcation may be simulated imposing a kinked VD. However, even though shifts in the axis of extension due to topography have been previously documented through analogue models (van Wyk de Vries and Merle, 1996), the performed experiments suggest that the lithospheric structures are too distant (low H/D ratio) to induce any deformation on the summit of Etna. Therefore, no significant direct effect of the activity of the ME can be expected to affect both the stability and volcanic activity in the upper part of the volcano. Set 1 experiments also show how the activity of a lithospheric boundary increments, on the long-term (>105 years), the gravitational gradient of the slope connecting the thicker continental lithosphere with the thinner oceanic lithosphere. This may thus provide an important contribution for the differential confinement at the base of the volcanic edifice, given by the topographic/bathymetric difference between the continental and oceanic crust. In summary, this experimental set suggests that, on the longer

term, the indirect topographic consequences of the activity of a lithospheric boundary may play a more important role on the instability of a volcano than any direct tectonic process. Experimental set 2 evaluated the possible role of a weak basement below the volcano on a mean term (b10 4 years). The overall behavior is similar to preexisting experiments and theoretical models (Borgia, 1994; Merle and Borgia, 1996; Borgia et al., 2005). The most interesting results of this set are: 1) whatever the boundary conditions (total or partial unbuttressing, cone elongation), radial graben accommodate the spreading of the cone above the weak basement; 2) peripheral strike-slip faults, radial graben on the lower slope and tangential folds at the volcano base develop after higher amounts of spreading, respectively. The radial graben is a consequence of the enlargement of the volcano base under a weak basement (Merle and Borgia, 1996). In fact, the weak basement enhances the radial spreading of the edifice, which increases its external circumference; in order to accommodate such an enlargement, radial graben forms. Similar tangential extension was observed also in analogue models simulating resurgence (Acocella et al., 2001). The radial growth of the edifice in turn induces compression at its base, in the weak basement, generating the tangential folds (Merle and Borgia, 1996). A differential unbuttressing (confinement) at the sides of the weak material may limit and focus the spreading process (Fig. 6d). Regardless of the initial eccentricity of the cone base (this varying between 1 and 0.8 in experiments 2b and 2c, respectively), the final edifice remains elongated parallel to the main spreading direction.

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Fig. 10. Summary of the experimental conditions (and the simulated time scale; left column), of the main results (central column), and their feasibility at Etna (right column).

Experimental set 3 showed how a pressurized intrusion may deform the overlying cone simulating periods of 101–102 years (medium-term behavior). The surface deformation pattern was significantly asymmetrical, with the main horizontal displacement on the eastern side (Fig. 8). This asymmetry is interpreted as the result of the lower confinement in the part of the basement of the cone where the scarp is located, which induces a local stress reorientation (e.g. Borgia et al., 1992: Froger et al., 2001; Solaro et al., 2010; Norini and Acocella, 2011). Experimental set 4 studied the surface deformation due to dike emplacement within the cone, within a simulating period of a very few days (short-term behavior). The direction of the dike, perpendicular to the basement slope, implies a topographic control responsible for the development of a minimum horizontal stress parallel to the basement slope, consistently with previous experiments (Fiske and Jackson, 1972). In addition, the resulting surface deformation pattern on the intruded cone was significantly asymmetrical, with the largest horizontal displacements towards the scarp side (see elaboration of kinematic data in Norini and Acocella, 2011). As the simulated dikes were centered below the cone summit, the asymmetric deformation is again interpreted as related to the differential buttressing, with minor confinement at the base of the cone where the scarp is located (Norini and Acocella, 2011). 5.2. Experimental constraints to the instability of Etna Set 1 experiment has shown how a lithospheric boundary may provide a steep slope and therefore a differential confinement at the base of the volcanic edifice. This feature matches the topographic and bathymetric configuration of the Etna area, with the mean altitude of the Appennine–Maghrebian Chain (1 km a.s.l., to the N and W) and the mean depth of the Ionian abyssal plain (− 2 km b.s.l., to the E). The gentler slope found in nature is attributed to the erosion

and sedimentation processes which have been occurring in the Sicilian offshore in the last few Ma (Argnani and Bonazzi, 2005, and references therein), which have not been simulated. The topographic and bathymetric configuration of the basement of Etna has been inferred to play a crucial role on the stability of the volcano (e.g. Froger et al., 2001), providing a differential confinement that decreases eastwards (Solaro et al., 2010). Alone, such a differential confinement appears insufficient to create any flank instability. However, as suggested by the experimental sets 2 to 4, the differential confinement at the volcano base may constitute a crucial preparing factor to enhance the deformation and the instability on one side of the volcano (Fig. 10). Without such a condition, any external triggering factor is expected to produce a symmetric (i.e. radial) deformation pattern (e.g. Acocella, 2005). Set 2 experiments are less straightforward to be applied to Etna. On the one hand, radial grabens at Etna have not been observed, making this lack the main obstacle to the application of the spreading model due to a weak basement (Norini and Acocella, 2011). Also, the upper edifice of Etna is N–S elongated, that is perpendicular to the main spreading direction. In the experiments, the elongation of the final cone is parallel to the main spreading direction, suggesting a limited effect of a weak basement on flank instability at Etna, on the upper edifice at least. On the other hand, compressional structures have been recently identified along a significant portion of the volcano base, through InSAR data (Solaro et al., 2010), and may be interpreted as resulting from spreading over a weak basement. The experiments suggest that these tangential compressional structures develop shortly after the onset of formation of the radial graben on the periphery of the volcano, at an advanced stage of spreading, corresponding to 10 2–10 3 years in nature. Therefore, to justify the presence of the tangential compressional structures, as well as the lack of the radial graben, which should form before, the following


possibilities have to be considered. a) The radial graben may currently be in an immature stage at Etna, making their identification difficult, also for any additional masking from volcanic activity. b) The peripheral extension associated with the radial graben in the experiments is being accommodated by the activity of a group of other structures predating the development of any radial graben at Etna; these structures may be largely reconciled with the main transtensive faults systems on the E flank, usually related to regional tectonics (including the Timpe Fault System, the onshore continuation of the ME) or with the transtensive structures induced by the collapse (including the Pernicana and Ragalna fault systems). The fact that recent and detailed InSAR data do not find evidence for the radial graben and, also, the dominant extensional component of motion shown by all the collapse structures on the flanks of Etna is ESE directed (Acocella et al., 2003; Solaro et al., 2010), suggest that hypothesis b) is more likely. Therefore, it is proposed that, while spreading due to a weak basement at Etna may be a viable, though limited, mechanism, its complete identification (highlighted by the formation of radial graben structures) is hindered by variously-oriented pre-existing structures, whose extensional component of motion replaces that which should have been accommodated by the radial graben (Fig. 10). In addition, the mean deformation rates of set 2 correspond to 101 cm/yr in nature, one order of magnitude larger than those deducible from InSAR data (Solaro et al., 2010). The discrepancy between experimental and natural rates implies two possibilities. One is that the experiments do not perfectly reproduce the natural case; an obvious limitation concerns the knowledge of the poorly constrained thickness and exact nature of the weak basement layer, which in the experiments is fixed to ~1 km. The discrepancy suggests that decreasing such a thickness would better scale the experimental rates to nature. Another possibility is that, if the thickness of the silicone proposed in the experiments is realistic, the viscosity of the weak basement as a whole may be one order of magnitude higher than previously inferred (that is, 1015 Pa s; Fig. 10; Palano et al., 2009). Finally, any spreading due to a weak basement at Etna may locally enhance magma intrusion and/or be activated by magma intrusion. Experimental sets 3 and 4 have shown an asymmetric deformation under the intrusion of magma analogues. The amount of deformation of the least buttressed flank, due to the change in slope below the summit, is several times that observed on the opposite flank (Figs. 8b and 9b). For comparison, available GPS data show that the areal extent of the eastern part of the volcano deformed by the inflation of the Mogi-like source between 1993 and 2001 is twice of that on the western side, whereas the amount of horizontal deformation is a very few times larger (Bonforte and Puglisi, 2006; Palano et al., 2008). Also, available data during the emplacement of the 2001 and 2002 dikes show that the horizontal displacement on the western side of the volcano was in the order of 10% with regard to that on the eastern side (Acocella and Neri, 2005; Solaro et al., 2010), similarly to what observed in the experiments. In terms of rates, the pressurization of a Mogi-like source in the experiments corresponds to deformation rates of 10 1–10 2 cm/yr in nature (Norini and Acocella, 2011), which are one order of magnitude larger with regard to the real case (Solaro et al., 2010). This discrepancy is probably related to the difficulty in properly modeling the poorly-known upper crustal structure below the volcano (Norini and Acocella, 2011) and in the limited knowledge of the mechanical parameters describing the block equilibrium (Battaglia et al., 2011). The emplacement of dikes results in a mean displacement rate of 10 0–10 1 cm/yr, which is in close agreement with field observations (Acocella and Neri, 2005; Palano et al., 2008; Solaro et al., 2010). More in general, the asymmetric deformation observed in the experiments can be related to the lower confinement in the eastern part of the volcano basement (Fig. 10; Froger et al., 2001; Solaro et al., 2010). Set 1 has shown how such an asymmetric deformation may result from the long-term activity of a lithospheric structure. The

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asymmetric deformation is responsible for the local reorientation of the minimum horizontal compressive stress, which becomes parallel to the basement slope, consistently with previous results (Fiske and Jackson, 1972; Acocella and Neri, 2009 and references therein). Even though the differential confinement itself is not sufficient to induce flank instability, under magmatic intrusions the differential confinement provides a consistent asymmetric deformation, which increases towards the basement slope. The importance of the differential buttressing, as evaluated in these experiments, provides a theoretical confirm for the crucial role of the topographic and bathymetric conditions along the volcano base in interpreting the recent deformation of Etna (Chiocci et al., 2011), as also revealed by InSAR data (Froger et al., 2001; Solaro et al., 2010). Also, while at Etna the weak basement provides a radial motion of the edifice, poorly sensitive to the differential buttressing, at least onshore (as testified by InSAR data; Solaro et al., 2010), the effect of the magmatic intrusions is particularly sensitive to the differential buttressing, amplifying the unstable area towards the least confined side (Fig. 10; Solaro et al., 2010). In summary, while the activity of the lithospheric boundary may indirectly provide the preparing conditions (differential unbuttressing) to induce the long-term (>105 years) collapse along the eastern flank of Etna, a weak basement (acting over a period b104 years) and, in particular, the intrusion of magma (101–2 years) may promote the destabilization of the volcano flanks on the shorter-term. In general terms, the experiments confirm that these processes are very effective in destabilizing the eastern and southern flank of the volcano (e.g. Puglisi et al., 2008), in a very similar fashion to what revealed by InSAR data (Solaro et al., 2010). However, this study has also shown that other secondary factors (including the importance of the extensional component of the pre-existing structures on the volcano slopes) may play an important role in better defining the instability at Etna (Chiocci et al., 2011). We expect that the example of the instability of the flanks of Mount Etna, as summarized here, may also enhance a better understanding of less studied or poorly known cases of volcano flank instabilities worldwide. 6. Conclusions The considered analogue models provide the following constraints: 1) The indirect topographic effect of the activity of a lithospheric boundary may play a more important role, on the long-term, on the instability at Etna than any direct tectonic process. In particular, the creation of differential buttressing conditions provides a crucial factor to prepare instability. 2) A weak basement may be a viable, though limited, mechanism to have flank instability at Etna, provided that pre-existing/regional structures with an extensional component replace the formation of any radial graben. 3) A pressurized Mogi-like source or dike emplacement may trigger and/or significantly enhance flank instability towards the least buttressed side. 4) More in general, the experiments confirm that the simulated processes are very effective in destabilizing the E flank of the volcano, in a very similar fashion to what recently revealed by InSAR data (Solaro et al., 2010). Acknowledgments Giuseppe Puglisi provided helpful discussions and Tiziana Apuani allowed a preliminary useful comparison with numerical models. Anna Nagay and Fabio Corbi helped during the analogue modeling at the Laboratory of Experimental Tectonics (LET), Roma Tre. Two anonymous reviewers and the Editor provided detailed and helpful comments, improving the work. This work was partially funded by INGV and the Italian DPC (DPC-INGV project V4 “Flank”).

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